Seasonal variability at the PIRATA sites
Among the four considered PIRATA buoy locations, which are representative for different regions within the ACT, large seasonal variability in the background setting is observed. Common to all regions within the ACT is the strong cooling of SSTs starting around April/May (Fig. 2a). The strongest cooling is observed on the equator at 10°W, which is sometimes referred to as the center of the ACT (Jouanno et al. 2011b). The structure of the seasonal variability of SSTs at the western (23°W) and eastern (0°E) edge of the ACT on the equator is very similar to 10°W. However, the southern ACT region exhibits a slower cooling phase followed by a delayed and more rapid warming phase. Towards the end of the year SSTs recover to warm levels at all locations reaching their maximum levels in March/April again.
The seasonal cycle of net surface heat fluxes from the TropFlux product (see Appendix 1 for a comparison between TropFlux and PIRATA) varies considerably within the ACT (Fig. 2b). The seasonal variation at all locations is mainly caused by variations in the incoming solar radiation and the latent heat flux, whereas the sensible heat flux and the outgoing long-wave radiation remain rather constant throughout the year (Foltz et al. 2003; Hummels et al. 2013). However, within the equatorial region, net surface heat flux is typically positive. In the western and central part of the equatorial ACT region (23°W and 10°W) the ML is warmed throughout the year by 50–100 W m−2. Note that here the net surface heat flux is calculated with the absorbed shortwave radiation, which is corrected for the amount of heat penetrating through the ML estimated from PIRATA sub-surface temperatures. Within the eastern equatorial ACT (0°E) the superposition of annual and semiannual cycles of the net surface heat fluxes leads to a net heat flux minimum from May to July when SSTs decrease. In the southern ACT region (10°S, 10°W) a strong annual cycle of net surface heat fluxes is observed, which includes a ML warming as well as a cooling phase. At this location the cooling phase of the ML due to a change in atmospheric forcing of nearly 200 W m−2 coincides with a decline of SSTs.
MLDs are in general shallow in the eastern equatorial Atlantic due to the shoaling thermocline. Accordingly, within the equatorial ACT region MLDs decrease from west (23°W) to east (0°E) (Fig. 2c). At all locations within the equatorial belt, MLDs exhibit a seasonal cycle with maximum MLDs of 40–50 m in boreal autumn and 15–20 m during boreal spring. Within the southern ACT region MLDs are generally larger than in the equatorial belt. In the south, the seasonal variation ranges from around 40 m in boreal winter/spring to maximum 80 m during boreal summer/autumn.
Winds in the ACT region are dominated by the westward trades (easterlies) surrounding the ITCZ, which migrates meridionally during the year. For the equatorial ACT region this migration imprints an annual signal on the wind stress magnitude with strongest winds in August, when the ITCZ is at its northernmost position (Fig. 2d). A weaker semi-annual signal, which peaks in April/May and September/October is superimposed on the dominant annual pattern yielding a double peaked structure. The southern ACT region exhibits only an annual variation in wind stress, which peaks in August, increasing the latent heat flux during this period. This increased latent heat flux contributes to the strong cooling via net surface heat fluxes in the southern ACT (Fig. 2b).
Zonal surface velocities in the equatorial ACT region, determined from a combination of ARGO float and surface drifter data, are subject to a relatively strong semi-annual cycle. This appears odd at first glance, as the wind forcing is dominated by the annual cycle with only a weak semi-annual imprint. However, the semi-annual cycle in zonal velocity was explained by the presence of a resonantly forced basin mode (Cane and Moore 1981; Thierry et al. 2004; Ding et al. 2009). The current dominating the equatorial region is the northern branch of the South Equatorial Current (nSEC) (Lumpkin and Garraffo 2005; Hummels et al. 2013), (Fig. 2e). During boreal summer maximum westward velocities are observed of about 0.25–0.5 m s−1 depending on the exact location within the equatorial belt. In the central and eastern equatorial Atlantic, zonal velocities even reverse in sign during the seasonal cycle. Due to the vanishing Coriolis parameter at the equator, meridional velocities might be directly forced by the meridional wind component. As winds at the equator have a southerly (from south to north) component and are largest in boreal summer and fall, meridional velocities are directed in the same direction (Rhein et al. 2010). This is also consistent with the results from Perez et al. (2013), where positive (northward) meridional velocities are found at the equator at 23°W and 10°W. However, their magnitude is significantly reduced compared to zonal velocities (Fig. 2f). In the southern ACT region zonal as well as meridional surface velocities are of reduced magnitude compared to the equatorial region and do not show a distinct seasonal variation. Surface velocities within this region are dominated by the Ekman flow. According to the steady trade winds (south-easterlies), the Ekman flow is directed towards the southwest throughout the year (Fig. 2e, f).
Turbulent mixing within the ACT
In this section, the new and unique data set of microstructure observations acquired in the central and eastern equatorial Atlantic (Hummels et al. 2013) is used to estimate the diapycnal ML heat loss directly, rather than relying on residual estimates of this quantity as has been done previously (Wang and McPhaden 1999; Foltz et al. 2003, 2013; Wade et al. 2011). The main findings relevant for this study from Hummels et al. (2013) concerning the regional and seasonal variability of turbulent mixing are summarized in the following: Turbulent dissipation rates (ɛ) at the equator are significantly increased in the upper thermocline compared to off-equatorial locations (cf. Fig. 3a). Within the equatorial region turbulent dissipation rates in the upper thermocline are:
-
1.
Elevated in the western equatorial ACT region in comparison to the eastern equatorial ACT region (cf. Figure 3a);
-
2.
Most intense mixing is observed in boreal summer in the whole equatorial ACT region;
Hummels et al. (2013) described a close correspondence between the seasonal and regional variability of background shear and stratification levels and ensemble mixing intensities, turbulent eddy diffusivities (K
ρ
) and diapycnal heat fluxes (J
h
): e.g. vertical shear squared of horizontal velocities (\(S^{2} = \left( {\partial U/\partial z} \right)^{2} + \left( {\partial V/\partial z} \right)^{2}\)) was found to reduce from 10°W towards the eastern equatorial ACT, while stratification (N2) increased. This reduces the likelihood of shear instabilities to occur in the Gulf of Guinea and indeed turbulent parameters such as dissipation rates as well as diapycnal ML heat losses were observed to decrease in magnitude from the western equatorial region towards the east. Despite stronger zonal subsurface velocities at 23°W compared to 10°W (Brandt et al. 2011a), shear levels are reduced at 23°W compared to 10°W (Jouanno et al. 2011b). This can be explained by the shoaling of the EUC towards the east, which limits the depth range of opposite flowing currents, westwards at the surface and eastwards in the subsurface. The highest shear levels in the central equatorial ACT region correspond to highest mixing activity. Within the southern ACT, shear levels were significantly reduced due to the lack of strong current features and turbulent mixing was found to be low. The relation between background shear and stratification conditions and turbulent mixing activity will be further analyzed later in this study.
According to the described variability in shear levels, average summer dissipation rates below the ML range from 3.5 × 10−7 W kg−1 in the eastern to 7 × 10−7 W kg−1 in the western equatorial ACT (Fig. 3a). In the southern ACT dissipation rates below the ML do not exceed 5 × 10−8 W kg−1. At a depth of 50 m below the ML, equatorial dissipation rates have dropped by about one order of magnitude (Fig. 3a). Inferred eddy diffusivities (Sect. 2.1.1) just below the ML range from 1.2 × 10−4 m2 s−1 in the eastern to 7 × 10−4 m2 s−1 in the western equatorial ACT, while values in the southern ACT only reach 1.6 × 10−5 m2 s−1 at maximum. The magnitude and vertical structure of the observed turbulent parameters agrees rather well with those inferred from microstructure measurement programs carried out in the central equatorial Pacific (Gregg et al. 1985; Peters et al. 1988; Moum et al. 1989; Lien et al. 1995). However, in comparison to the central equatorial Pacific, Hummels et al. (2013) report for the equatorial ACT region a reduction in the night time enhancement of turbulence, which is referred to as deep cycle turbulence (Moum and Caldwell 1985). Nevertheless, to avoid possible biases due to an unevenly distributed sampling time of profiles during the day, mean profiles of turbulent parameters are derived here by separately averaging measurements collected during the day (08:00–20:00) and night (20.00–08:00) before calculating mean profiles. The average of the mean day and mean night profiles are then further incorporated into the ML heat budgets.
In order to obtain the turbulent contribution of diapycnal heat fluxes to the ML heat budget, the transition zone between the base of the ML and the stratified region below has to be accurately resolved. In the equatorial Atlantic as well as in the Pacific, profiles of the diapycnal heat flux are highly divergent below the ML. Maximum values are found at the base of the ML that rapidly decrease in deeper layers (Fig. 3d; Lien et al. 2008). In most of the mean profiles (Fig. 3d), diapycnal heat flux at 20 m below the ML is significantly reduced. Hence, the amount of heat being extracted from the ML into the interior ocean is characterized by the diapycnal heat flux in a rather narrow layer. Here, diapycnal ML heat loss is determined by averaging the diapycnal heat flux profiles in the interval MLD + 5 m to MLD + 15 m. The reason to use this averaging interval is to ensure values from within the ML to be excluded from the estimate as the method from Osborn (1980), which we use here, is only valid in stratified sheared flow. Due to strong variability in stratification within a few meters below the ML, and due to the fact that stratification for the Osborn parameterization needs to be calculated over a larger length scale than turbulent overturns (that can be as much as several meters due to the strong mixing there), we decided to use this depth interval. Due to the elevated vertical divergence of the heat flux profiles, this approach leads to estimates of diapycnal heat loss of the ML being biased low. The error of this approach can be estimated from the average heat flux profiles by extrapolating from MLD + 10 m to the MLD. The extrapolation needs to be performed as the diapycnal heat flux obtained directly below the ML with the Osborn method can not be considered reliable as explained above. Overall, the diapycnal heat flux reduces by about 30 % at MLD + 10 m compared to the value directly below the ML.
Note that the MLD from vertically high resolved CTD profiles was generally calculated using the temperature threshold criterion with ΔT = 0.2 °C compared to the SST. Heat flux profiles and inferred diapycnal ML heat losses were calculated separately for every station. Subsequently, these station averaged profiles of the different sections of an individual cruise between 2°S and 1.5°N were averaged in day and night ensemble (see above) to obtain a single estimate that was taken as representative for diapycnal heat loss of the ML for the month in which the measurements were collected. Uncertainties for each individual estimate were calculated using Gaussian error propagation as described in Hummels et al. (2013).
Mixed layer heat budget
In the following, the contributions to the ML heat budget derived from PIRATA observations, climatological data sets as well as microstructure observations are combined at the four different locations within the ACT. Note that the individual contributions to the ML heat budget at 0°N, 10°W were already analyzed in Hummels et al. (2013). However, to achieve a consistent comparison between heat budgets at the different PIRATA locations, the analysis of the heat budget at 0°N, 10°W is repeated here partly using different data products (concerning the surface velocities and the net surface heat fluxes) compared to Hummels et al. (2013).
Several modeling as well as observational studies have addressed the ML heat budget in the eastern equatorial Atlantic previously (Foltz et al. 2003; Peter et al. 2006; Jouanno et al. 2011b; Wade et al. 2011; Hummels et al. 2013). Among these studies, the definition of high and low frequency advection terms, referred to as mean and eddy advection here (Sect. 2.4), varied. Hence, in order to compare the results amongst the different studies it has to be clarified which oceanic processes are attributed to the different terms. As has been pointed out in the tropical Pacific analysis by Wang and McPhaden (1999) processes reflected in the eddy advection term close to the equator, particularly in the meridional component, significantly depend on the latitudinal averaging interval over which the heat budget analysis is performed. For local heat budgets or budgets performed for small regions around the equator (i.e. ±2° in latitude), the eddy advection warms the ML as the effect of TIWs laterally advecting warm waters are explicitly resolved (Foltz et al. 2003; Jochum and Murtugudde 2006; Peter et al. 2006). When budgets are performed over a larger meridional extent, eddy advection will cool the ML as TIW contributions are averaged out and the Ekman divergence dominates. This study focuses on local heat budgets at the locations of the PIRATA buoys and thus requires evaluating the eddy heat fluxes as local as possible. Opposing this minimum regional extent requirement is the accuracy of velocity and SST gradient data, for which the statistical reliability of the individual variables increases when they are averaged over larger meridional intervals. As a compromise between locality and statistical reliability, a 2° latitudinal and longitudinal interval is used to evaluate mean and eddy advection terms. In doing so, the warming effect of TIWs will dominate the eddy advection term.
When the net surface heat flux is compared to the observed heat storage, large negative residuals are evident in the western (23°W) and central (10°W) equatorial ACT and a reduced residual in the eastern (0°E) equatorial ACT (Figs. 4, 5). In contrast, within the southern ACT region, SST variability during ACT development can virtually be explained by the variability in net atmospheric forcing. Hence, the residuals in the equatorial region need to be explained by different oceanic processes, which probably also vary in their relative contribution within the seasonal cycle.
In the following, the respective contributions of atmosphere and ocean processes to the warming and cooling of the ML during the absence, development and mature phase of the ACT will be discussed. As the focus of this study is on the seasonal variability of the individual contributions to the ML heat budget, annual and semi-annual harmonics were fitted to the individual terms (except for the diapycnal ML heat loss) before illustration (Fig. 5).
0°N, 23°W
At the western edge of the ACT (23°W) in the central equatorial Atlantic, the ML is warmed by net atmospheric forcing (Fig. 2b, 5a) and by eddy heat advection. In the central equatorial Atlantic eddy heat advection is predominately controlled by TIWs (Wang and McPhaden 1999; Jochum and Murtugudde 2006; Peter et al. 2006). TIW activity here was reported enhanced in the beginning of the year, in boreal summer and autumn (Bunge et al. 2007; von Schuckmann et al. 2008), which agrees with periods of elevated eddy advection in this analysis (Fig. 5a). Cooling of the ML is achieved by subsurface processes (diapycnal mixing and entrainment) as well as mean heat advection (Fig. 5a).
During the absence of the ACT (January to April), the ML balance is dominated by net atmospheric forcing and eddy heat advection, the latter contributing to a warming of up to 50 W m−2 in January. During March and April, when the tropical Atlantic is uniformly warm and the meridional gradient of SST is very weak, the eddy heat advection reduces to zero.
During the development phase of the ACT (May to August) the net surface heat flux increases, mainly due to increased incoming solar radiation, which counteracts the observed cooling of SSTs. Eddy advection dominated by the effect of TIWs as discussed above contributes to the warming of the ML. However, there is a strong increase in ML cooling from zonal heat advection (60 W m−2) and entrainment (25 W m−2) during this phase. The elevated zonal advection term is due to the persistent westward flow (Fig. 2e) advecting cooler surface waters from the central ACT towards 23°W. In addition, elevated diapycnal heat loss at and below the ML further contributes to cool the ML: the two independent estimates for June from 2006 and 2011, 58 W m−2 (M68/2) and 54 W m−2 (MSM18/2) respectively, agree well. The cooling dominated by the diapycnal heat flux and zonal heat advection is strong enough to reduce SSTs despite the warming due to net surface heat fluxes and eddy advection.
Within the mature phase of the ACT (August to the end of the year) the net surface heat flux significantly warms the ML. Eddy advection still contributes significantly (50 W m−2) to the warming. Cooling provided by zonal heat advection decreases at the beginning of the mature phase, due to the reduction in surface velocities associated with the nSEC (Fig. 2e). Towards the end of the year zonal heat advection re-intensifies in accordance with re-intensified zonal velocities associated with the nSEC. The diapycnal heat flux during this period is still large. The two independent estimates for November from 2009 and 2012 yielded 45 W m−2 (MSM22) and 30 W m−2 (M80/1) respectively. The meridional heat advection increases within this phase due to increasing meridional velocities (Fig. 2f) and increasing MLD (Fig. 2c). Entrainment still contributes a cooling of about 15 W m−2 during this phase. However, the gradual overall reduction of the cooling terms lead to a gradual increase in SSTs towards the end of the year.
Comparison of the sum of the individual terms contributing to the ML heat budget and the observed heat storage reveals a large residual of 30–100 W m−2 throughout the year, if the contribution of the diapycnal heat flux is omitted (Fig. 5b). This was also reported in the recent study of Foltz et al. (2013). Similar residuals (around 80 W m−2) have been reported in previous observational studies from the central equatorial Pacific (Wang and McPhaden 1999) as well as from the western equatorial ACT (Wade et al. 2011), where the diapycnal contribution could not be estimated. Including the resolved seasonal variability of the diapycnal heat flux into the sum of terms reduces the residual in boreal summer and November by more than a factor of 2 and closes the heat budget at least within the uncertainties (Fig. 5b). The diapycnal ML heat loss together with mean zonal advection are identified as the dominant contribution to the cooling of SSTs during ACT development at 0°N, 23°W.
0°N, 10°W
In the center of the ACT (10°W) the ML is warmed by the atmosphere and eddy advection as was observed for 23°W. The ML is cooled by subsurface processes (diapycnal mixing and entrainment) as well as the meridional heat advection (Fig. 5c). Zonal heat advection at this site is significantly reduced in magnitude compared to the western edge of the ACT (23°W). Although westward surface velocities associated with the nSEC are also intensified at this location (Fig. 2e), the low zonal temperature gradient in the center of the ACT and shallow MLDs lead to reduced zonal heat advection compared to 23°W.
During the absence of the ACT (January to April) the ML heat budget at 10°W is dominated by net atmospheric forcing. The largest oceanic contribution is the meridional heat advection cooling the ML, which balances the warming via zonal heat advection and eddy advection. The meridional velocity is as explained previously a direct response to the meridional wind forcing. Although elevated ML cooling due to diapycnal heat fluxes is anticipated during this period, this study lacks observational support for this hypothesis.
Within the development phase of the ACT (May to August), the net heat flux from the atmosphere increases due to the reduction in latent heat flux due to reduced wind speed and an increase in the incoming solar radiation (Figs. 2b, 5c). Additional warming is provided by the eddy advection due to TIWs, similar as discussed for 23°W. However, strong subsurface cooling leads, despite these warming terms, to a cooling of the ML. The relative constitution of the subsurface cooling at this location differs from the observations at 23°W: Zonal heat advection is significantly reduced compared to 23°W, whereas entrainment is of similar magnitude cooling the ML at a rate of about 20 W m−2, which agrees with the results obtained by Rhein et al. (2010). The striking difference is the clear dominance of the diapycnal ML heat loss of up to 90 W m−2 over the other oceanic cooling contributions. As detailed in Hummels et al. (2013) elevated vertical shear of horizontal velocities increases the occurrence of shear instabilities leading to the elevated diapycnal heat fluxes during this phase.
During the mature phase of the ACT (August towards the end of the year) atmospheric warming stays on a rather high level. The cooling due to the diapycnal heat flux is still of considerable magnitude, but decreasing. The dominance of the diapycnal heat flux within the subsurface cooling terms reduces in favor of the meridional heat advection, which reaches 60 W m−2 towards the end of the year. This increase in meridional heat advection is due to enhanced meridional velocities presumably caused by increased southerly winds, temperature gradients and MLDs (Fig. 2). The warming effect of the eddy advection reduces during the mature phase and is accompanied by the zonal heat advection actually warming the ML towards the end of the year. This is due to the changing sign of the zonal velocity (the reversal of the nSEC between August and October; Fig. 2e) and the subsequent sign change for the zonal temperature gradient occurring in November and December (not shown). The gradual reduction in the total subsurface cooling leads to a gradual increase in SSTs during this phase.
Comparison of the sum of the individual terms of the ML heat budget to the observed heat storage reveals a large residual of up to 110 W m−2 when omitting the contribution of the diapycnal heat flux during boreal summer and autumn (Fig. 5d). This was already reported by Foltz et al. (2003), who performed a similar study at this location. Consideration of the diapycnal heat flux as a contributing term yields in a closure of the budget within the uncertainties from June to November. However, as further described below, ML cooling due to diapycnal mixing is likely to contribute to the ML budget also during the absence of the ACT.
0°N, 0°E
At the eastern edge of the ACT (0°E) the ML is mainly warmed by the net atmospheric forcing and cooled by the diapycnal heat flux (Fig. 5e). The other oceanic contributions do not exceed 20 W m−2 throughout the year. Zonal heat advection at this location is negligible throughout the year due to reduced zonal temperature gradients and MLDs (Figs. 1, 2c), whereas entrainment acts to slightly cool the ML throughout the year, similar to what has been observed at 10°W. However, as pointed out above the net surface heat flux is significantly reduced in the eastern, equatorial ACT region compared to the more western and central locations (23°W, 10°W, Figs. 2b, 4). Accordingly, less subsurface cooling is required to decrease SSTs at this location.
During the absence of the ACT (January to April) in the beginning of the year the ML heat budget is determined by net atmospheric forcing and oceanic contributions do not exceed 10 W m−2.
Incoming solar radiation reduces during ACT development (May to August) while the latent heat flux slightly increases, which leads to the reduction in net surface heat fluxes during this period. Concurrently, the diapycnal heat flux as inferred from microstructure observations during June 2006 and 2007 increases to its maximum value at this location of 21 and 29 W m−2 respectively and dominates the subsurface cooling (Fig. 5e). Note that at this location the seasonal variability of the diapycnal heat flux was composed from available estimates at 0°N, 0°E as well as around 0°N, 2°E. MLDs are observed extremely shallow at this longitude (Fig. 2c). Hence, the additional subsurface cooling by diapycnal mixing, which superimposes on the reduced net surface heat flux, is sufficient for the strong decrease in SSTs. Eddy advection is negligible during this phase. At this longitude within the Gulf of Guinea TIWs have not been detected. Instead, Athie and Marin (2008) as well as the numerical simulation of Jouanno et al. (2013) suggest intraseasonal variability in the Gulf of Guinea dominated by wind-forced Yanai waves having long zonal wavelengths and a period between 10 and 20 days. From our analysis it seems that the effect of these waves with long zonal wavelengths on the eddy advection is significantly reduced compared to the effect of TIWs in the western equatorial ACT. Nevertheless, they might contribute to enhance vertical shear of horizontal velocity and thereby favor diapycnal mixing at this location as suggested by Jouanno et al. (2013).
During the mature phase of the ACT (August to the end of the year), a similar evolution towards the end of the year is observed as at 10°W on the equator. Net surface heat fluxes increase due to the increase in the incoming solar radiation. Meridional and eddy heat advection both contribute about 20 W m−2 during this phase but with opposite sign. The cooling by the diapycnal heat flux reduces, which leads together with the increased warming by the atmosphere, to the retraction of the ACT towards the end of the year.
Comparison of the sum of individual terms to the ML heat budget and the observed heat storage reveals a residual of up to 30 W m−2 when omitting the contribution of the diapycnal heat flux (Fig. 5f), which is within the uncertainties at this location. However, if the diapycnal heat flux is included, the residual reduces minimum by a factor of 2 (Fig. 5f). Even if reduced in magnitude compared to the western and central equatorial region, the diapycnal heat flux provides the largest subsurface cooling term within the development phase of the ACT and hence seems to supply the essential contribution to cool SSTs.
No detailed observational study of the ML heat budget as far east as 0°E in the Gulf of Guinea has been published so far. Wade et al. (2011) determined the individual terms of the ML heat budget for considerable larger regions (about 5° latitude and 9° longitude). Their box 5, representative for a region including 0°N, 0°E, shows a considerable larger residual term of maximum 80 W m−2 during ACT development, which they associated with the diapycnal heat flux. Note, though that their box 5 extends until 6°W, where the diapycnal heat flux may still be elevated compared to 0°E. In addition, their estimate of the net surface heat flux within this box ranges from 50 to 120 W m−2 and is above our estimates at 0°N, 0°E especially during ACT development (Fig. 2b). Thus, for their study a larger cooling by oceanic processes is required to match the observed heat storage.
10°S, 10°W
Investigation of the background setting at the different PIRATA sites already revealed fundamental differences between the equatorial and the southern ACT region (Fig. 2). The net surface heat flux actually cools the ML during ACT development in contrast to the equatorial ACT region, where net atmospheric warming is observed throughout the year and at all locations. In general, surface velocities are weaker at this location compared to the equatorial region (Fig. 2e, f) and are not associated with distinct current features, but with the Ekman flow. Meridional heat advection tends to warm the ML at this location, while entrainment, zonal advection and eddy advection mostly cool the ML.
During the absence of the ACT (January to April) the increase in ML heat storage is balanced by net atmospheric fluxes, warming the ML (Fig. 5g).
During the development of the ACT (May and August) net surface heat fluxes significantly cool the ML by up to 90 W m−2 (Figs. 2b, 5g) due to the increase of the latent heat flux associated with increased winds as well as a reduction in the incoming solar radiation (Figs. 2d, 5g). Surface currents associated with the Ekman flow are directed towards the southwest. Meridional advection contributes to warm the ML during this period due to the southward advection of warmer waters from the north (see Fig. 1). The other advection terms as well as entrainment are rather small during the development phase of the ACT. Zonal heat advection is small due to the lack of a significant zonal temperature gradient (not shown). The diapycnal heat flux inferred from microstructure observations during June 2006 provides a negligible contribution to the cooling (Fig. 5g). Thus, the cooling of the ML and hence SSTs during ACT development at this location is dominated by the net atmospheric forcing. Hence, the deepening of the ML from April to August can not be associated with mechanical mixing, but must be due to an increase in wind stress curl or surface buoyancy flux.
During the mature phase of the ACT (August to the end of the year) the net surface heat flux increases and gradually warms the ML again. This is due to the increase in incoming solar radiation and a reduction of the latent heat flux due to decreasing winds. The diapycnal heat flux has not been estimated at this location during this phase in the seasonal cycle. However, highest mixing activity throughout the equatorial ACT was observed in boreal summer (Hummels et al. 2013), when vertical shear of horizontal velocities is strong. As the vertical shear at this location is in general rather low due to the lack of strong current features, it seems unlikely that this term gives an important contribution to the ML heat budget during this phase.
The sum of contributing terms closely follows the evolution of the observed heat storage at this location (Fig. 5h). The residual is negligible throughout the year in agreement with the previous estimates of Foltz et al. (2003). This agrees with the fact that the diapycnal heat flux at this location was estimated rather low during the main mixing season of the equatorial ACT (Figs. 3, 5g). The decrease in SSTs at this location during ACT development is governed by the net atmospheric forcing.